cloud system resolving model study of the roles of deep ...gated in a 7-day 3-d 248×248km2...

Atmos. Chem. Phys., 8, 2741–2757, 2008www.atmos-chem-phys.net/8/2741/2008/© Author(s) 2008. This work is distributed underthe Creative Commons Attribution 3.0 License.

AtmosphericChemistry

and Physics

Cloud system resolving model study of the roles of deep convectionfor photo-chemistry in the TOGA COARE/CEPEX region

M. Salzmann1, M. G. Lawrence1, V. T. J. Phillips2, and L. J. Donner3

1Max-Planck-Institute for Chemistry, Department of Atmospheric Chemistry, PO Box 3060, 55020 Mainz, Germany2Department of Meteorology, University of Hawaii at Manoa, 2525 Correa Road, Honolulu, HI 96822, USA3Geophysical Fluid Dynamics Laboratory, NOAA, Princeton University, PO Box 308, Princeton, NJ 08542, USA

Received: 30 October 2007 – Published in Atmos. Chem. Phys. Discuss.: 10 January 2008Revised: 7 May 2008 – Accepted: 7 May 2008 – Published: 22 May 2008

Abstract. A cloud system resolving model including photo-chemistry (CSRMC) has been developed based on a proto-type version of the Weather Research and Forecasting (WRF)model and is used to study influences of deep convectionon chemistry in the TOGA COARE/CEPEX region. Lat-eral boundary conditions for trace gases are prescribed fromglobal chemistry-transport simulations, and the vertical ad-vection of trace gases by large scale dynamics, which is notreproduced in a limited area cloud system resolving model, istaken into account. The influences of deep convective trans-port and of lightning on NOx, O3, and HOx(=HO2+OH),in the vicinity of the deep convective systems are investi-gated in a 7-day 3-D 248×248 km2 horizontal domain sim-ulation and several 2-D sensitivity runs with a 500 km hor-izontal domain. Mid-tropospheric entrainment is more im-portant on average for the upward transport of O3 in the 3-Drun than in the 2-D runs, but at the same time undiluted O3-poor air from the marine boundary layer reaches the uppertroposphere more frequently in the 3-D run than in the 2-Druns, indicating the presence of undiluted convective cores.In all runs, in situ lightning is found to have only minor im-pacts on the local O3 budget. Near zero O3 volume mixingratios due to the reaction with lightning-produced NO areonly simulated in a 2-D sensitivity run with an extremelyhigh number of NO molecules per flash, which is outsidethe range of current estimates. The fraction of NOx chemi-cally lost within the domain varies between 20 and 24% inthe 2-D runs, but is negligible in the 3-D run, in agreementwith a lower average NOx concentration in the 3-D run de-spite a greater number of flashes. Stratosphere to troposphere

Correspondence to:M. Salzmann([email protected])

transport of O3 is simulated to occur episodically in thin fil-aments in the 2-D runs, but on average net upward transportof O3 from below∼16 km is simulated in association withmean large scale ascent in the region. Ozone profiles in theTOGA COARE/CEPEX region are suggested to be stronglyinfluenced by the intra-seasonal (Madden-Julian) oscillation.

1 Introduction

Tropospheric ozone is an important greenhouse gas and a keyplayer in tropospheric chemistry. Over the remote tropicaloceans deep convection rapidly moves O3-poor air from themarine boundary layer (MBL) to the upper troposphere (UT)(e.g.Lelieveld and Crutzen, 1994; Wang et al., 1995; Kleyet al., 1996; Pickering et al., 2001) while over polluted con-tinental areas the transport of ozone and its precursors playsan important role for the upper tropospheric ozone budget(e.g.Dickerson et al., 1987; Pickering et al., 1990; Lawrenceet al., 2003). Deep convection can also bring O3-rich air fromthe lower stratosphere to the troposphere (e.g.Wang et al.,1995; Poulida et al., 1996; Stenchikov et al., 1996; Dye et al.,2000; Wang and Prinn, 2000) and air originating in the mid-troposphere is carried upwards in association with mid-levelinflow into tropical squall lines (Scala et al., 1990) or mid-latitude storms (Wang and Chang, 1993). Furthermore, light-ning NOx can enhance O3 production in deep convective out-flow hundreds of kilometers downwind of deep convection(e.g.Pickering et al., 1990, 1993, 1996). In the troposphere,downward transport takes place in association with rear in-flow into mesoscale convective systems (e.g.Scala et al.,1990; Salzmann et al., 2004). WhileKley et al.(1997) arguedthat observed sharp ozone gradients indicate a lack of vertical

Published by Copernicus Publications on behalf of the European Geosciences Union.

http://creativecommons.org/licenses/by/3.0/

2742 M. Salzmann et al.: Deep convection and chemistry in the TOGA COARE/CEPEX region

mixing across the tropopause, dry layers with high ozonemixing ratios are common in parts of the marine tropical tro-posphere and have been suggested to originate either fromdeep convection over continents lifting polluted air or fromthe stratosphere (Newell et al., 1996). In global chemistry-transport models, deep convective transport, which is typi-cally parameterized using simplistic mass flux approaches,constitutes a major source of uncertainty (Mahowald et al.,1995; Lawrence and Rasch, 2005).

Wang and Prinn(2000) studied a CEPEX storm (CEPEX:Central Equatorial Pacific Experiment;Ramanathan et al.,1993) using 2-D and 3-D cloud resolving model simulationsincluding chemistry and found that downward transport ofO3-rich air in association with deep convection, which tookplace in thin filaments in their simulations, lead to increasesin upper tropospheric O3 mixing ratios in spite of the upwardtransport of O3-poor air. They suggested O3 loss due the re-action with lightning-produced NO in cloud anvil regions canhelp explain the occasionally observed layers with extremelylow concentrations of O3 in the upper troposphere. Further-more, they found that a considerable fraction of lightningNOx was photochemically converted to HNO3 in the vicinityof the storms. This fraction decreased with increasing NOproduction from 28% in a 2-D sensitivity run with a fairlylarge NO production to 12% in a 2-D run with an extremelylarge NO production. At very high simulated NOx concen-trations they found NOx loss to be HOx limited. Althoughlightning is much less frequent over oceans than over land(e.g.Christian et al., 2003), lightning NO constitutes one ofthe most important sources of NOx over the remote tropi-cal oceans next to ship emissions and the thermal decom-position of peracetyl nitrate (PAN), which is advected overlong distances at high altitudes and subsequently transporteddownwards. Large photochemical losses of lightning NOxin the vicinity of the storms would therefore be important inspite of the large uncertainty still related to the amount of NOproduced per flash. Globally, lightning has been shown tohave an important impact on OH and the lifetime of methane(Labrador et al., 2004; Wild, 2007).

In this study we use a so-called cloud system resolvingmodel setup including chemistry (CSRMC) in order to studythe impact of multiple storms and mesoscale convective sys-tems over a period of several days. Horizontal advectionin and out of the domain is taken into account by specify-ing lateral boundary conditions for trace gases. Furthermore,the mean ascent in the domain (associated with the Hadleyand Walker cell and the intra-seasonal oscillation), which isnot reproduced by limited area cloud system resolving mod-els, is taken into account as inSalzmann et al.(2004) basedon mean vertical velocities derived from observations overthe tropical West Pacific east of Papua New Guinea duringTOGA COARE (Webster and Lukas, 1992; Ciesielski et al.,2003). Taking into account the mean ascent largely com-pensates the mass balancing mesoscale subsidence betweenclouds which would otherwise erroneously be forced to take

place in the domain (Salzmann et al., 2004) and for examplelead to a significant overestimate of the thermal decomposi-tion of PAN. We focus on the following issues:

1. The influence of NO production by lightning on NOx,O3, and HOx, and the sensitivity of the model results tothe dimensionality (2-D or 3-D) of the model.

2. Deep convective ozone transport (especially with regardto implications for its treatment in global chemistry-transport models), and, closely related to this, the roleof regional scale ozone transport and locally producedlightning NOx in explaining ozone profiles in the TOGACOARE/CEPEX region.

We also discuss the ozone results in the light of very lowupper tropospheric ozone concentrations observed duringCEPEX (Kley et al., 1996, 1997). Unfortunately, chemistryand flash rate observations are not available for the episode(19–26 December 1992) simulated in this model sensitivitystudy. In order to constrain the simulated flash rates and forthe discussion of our chemistry results, we instead rely onobservations from nearby regions and/or at other times.

The CSRMC is described in the next section. In Sects.3–5 results for NOx, O3, and HOx are presented. The influ-ence of the intra-seasonal (Madden Julian) oscillation (ISO)on ozone profiles is discussed in Sect.6 based on output fromthe global Model of Atmospheric Transport and Chemistry -Max Planck Institute for Chemistry version (MATCH-MPIC,Lawrence et al., 1999b; von Kuhlmann et al., 2003, and refer-ences therein). The CSRMC simulated meteorology (similarto Salzmann et al., 2004) and details of the CSRMC suchas the algorithms used for identifying updrafts and anvilsin the lightning parameterization are described in the elec-tronic supplementhttp://www.atmos-chem-phys.net/8/2741/2008/acp-8-2741-2008-supplement.pdf. Additional materialproviding a rather comprehensive view of the details of oursimulations and a brief description of MATCH-MPIC is alsofound in the supplement.

2 Model description and setup

The meteorological component of the cloud system re-solving model including photo-chemistry (CSRMC) isbased on a modified height coordinate prototype ver-sion of the Weather Research and Forecasting (WRF)model (Skamarock et al., 2001). Modifications to thebase model and the meteorological setup are largely iden-tical to Salzmann et al.(2004) and are described inthe electronic supplementhttp://www.atmos-chem-phys.net/8/2741/2008/acp-8-2741-2008-supplement.pdf(Sect. S1.1and Sect. S1.2). A “background” CH4−CO−HOx−NOx tro-pospheric chemistry mechanism with additional reactions in-volving PAN (peroxy acetyl nitrate, CH3C(O)O2NO2), andloss reactions of acetone (CH3COCH3) which is based on

Atmos. Chem. Phys., 8, 2741–2757, 2008 www.atmos-chem-phys.net/8/2741/2008/

http://www.atmos-chem-phys.net/8/2741/2008/acp-8-2741-2008-supplement.pdf




M. Salzmann et al.: Deep convection and chemistry in the TOGA COARE/CEPEX region 2743

the mechanism from MATCH-MPIC (von Kuhlmann et al.,2003) is used in the present study (Sect. S1.3). Photoly-sis rates are computed using the computationally efficientscheme byLandgraf and Crutzen(1998). A simple pa-rameterization of the production of nitrogen oxide by light-ning (see below) and a scheme for dry deposition over wa-ter (based onGanzeveld and Lelieveld, 1995, and referencestherein) have been implemented in the CSRMC. For solubletrace gases the uptake into or onto, release from, transport to-gether with, and mass transfer between different model cate-gories of hydrometeors (cloud droplets, rain, small ice parti-cles, graupel, and snow) are calculated as inSalzmann et al.(2007), see also Sect. S1.5. For HNO3 and H2O2 ice up-take from the gas phase is taken into account (Sect. S1.6).Aqueous phase reactions are currently not explicitly includedin the CSRMC, except for the hydrolysis reaction of N2O5.Dissociation reactions are, however, to some extent implic-itly included via the use of “effective” Henry’s law coeffi-cients in the mass transfer equations (Sect. S1.5). The rate ofchange of the gas phase concentration of thei-th trace gas is:

∂tCi,g= − ∇ ·(vCi,g

)+ ∂tCi,g|vlsat + ∂tCi,g|lno

+∂tCi,g|turb + ∂tCi,g|solu + ∂tCi,g|drydep

+∂tCi,g|chem (1)

where ∂t is the partial derivative with respect to time,v=(u, v,w) is the three dimensional wind vector,∂tCi,g|vlsatis the rate of change due to the vertical advection by the largescale vertical motion which is not reproduced by the CSRMC(Salzmann et al., 2004, see also Sect. S1.2),∂tCi,g|lno is therate of NO production by lightning,∂tCi,g|turb is the rateof change due to turbulent mixing by sub-grid scale eddies(parameterized applying Smagorinsky’s closure scheme,e.g. Takemi and Rotunno, 2003), ∂tCi,g|solu is the uptakeor release rate of soluble gases by or from hydrometeorsin the liquid or ice phase (Sect. S1.6),∂tCi,g|drydep is thedry deposition rate, and∂tCi,g|chem is the rate of changedue to gas phase chemical reactions. The advection of tracegases, of water vapor, and of hydrometeors in the liquid andice phase is calculated using the monotonic algorithm byWalcek(2000).

Lateral boundary conditions for trace gases are derivedfrom 3-hourly MATCH-MPIC output (see Sect. S2 andSect. S8). In the 3-D run (LTN3D), the horizontal domainsize is 248×248 km2 and the horizontal resolution is 2 km.The vertical resolution is 500 m in the troposphere with in-creasing grid spacings above 19 km and a total of 46 levelsup to 24707 m. A timestep of 8 s is used. In the 2-D runs(NOLTN, LTN1, LTN2, LTNWP, and LTNHWP), the hor-izontal domain extension is 500 km and the domain is ori-ented in East-West direction. The horizontal resolution isagain 2 km and the vertical resolution is 350 m in the tropo-sphere with a timestep of 5 s. The meteorological setup inthe 2-D runs is identical toSalzmann et al.(2004), while the

timestep and the vertical grid spacing were increased for the3-D run. An overview of the meteorological results is pre-sented in Sect. S3.

2.1 Simple lightning NO parameterization

Lightning NO production is parameterized as follows: As afirst step, updrafts are localized by fitting rectangles in thehorizontal plane enclosing areas where the vertical veloc-ity exceeds 5 m s−1. This is achieved by recursively cuttinglarger rectangles into smaller rectangles until no new edgesare found withwmax≤5 m s−1 everywhere (for a more de-tailed description of the algorithm see Sect S1.4). For eachupdraft, a flash rateFPR (in flashes per minute) is calculatedfrom:

FPR=α · 5 · 10−6wkmax. (2)

wherewmax is the maximum vertical velocity in meters persecond. Based on empirical relationships,Price and Rind(1992) derived an exponent ofk=4.54, which is somewhatless than the model study based estimate of at least 6 byBaker et al.(1995). Pickering et al.(1998), who amongother convective systems studied a squall line during TOGACOARE, used different values for different model simulatedstorms in order to obtain results comparable to observed flashrates. For their 2-D TOGA COARE simulation, they in-creased the exponent from 4.54 to 5.3.α is an empiricalscaling factor which is adjusted to improve the agreementwith available flash rate observations. For a mid-latitude su-percell stormFehr et al.(2004) applied a scaling factor of0.26. Here, sensitivity studies using various exponents andscaling factors were conducted (Table1, Sect.3).

The ratioZ=NIC/NCG of the number of intra cloud (IC)flashes to the number of cloud to ground (CG) flashes caneither be prescribed in the CSRMC or estimated from theempirical relationship (Price and Rind, 1993):

Z= 0.021(1h)4− 0.648(1h)3

+ 7.493(1h)2

− 36.54(1h) + 63.09 (3)

for clouds with1h≥5.5 km, where1h=zctw−z0◦C is thevertical distance between the 0◦C isotherm and the cloudtop at the location of the maximum vertical updraft velocity(zctw). The cloud top heightzct is defined for each verticalcolumn as the level above which the sum of the mass mix-ing ratios of all hydrometeor categoriesqtotm=qcloud water+

qrain + qice + qgraupel+ qsnow drops to below 0.01 g kg−1 inall overlying levels. For clouds with1h<5.5 km all flashesare assumed to be IC flashes based on the observation thatvery little lightning is produced by clouds with1h<5.5 kmand that clouds produce almost exclusively IC flashes dur-ing their growth stage before they start producing CG flashes(Price et al., 1997; Pickering et al., 1998).

The vertical distribution of lightning channel segmentsis essentially that ofDeCaria et al.(2000, 2005), i.e., CG

www.atmos-chem-phys.net/8/2741/2008/ Atmos. Chem. Phys., 8, 2741–2757, 2008


Table 1. Lightning NO sensitivity runs.

Label molec molec α in k in NCGb NIC

b Zc Ptotd Ptot

d Ptotd

CGa ICa Eq. (2) Eq. (2) CG IC

LTN3D 10 5 0.06 4.5 1021 10644 10.43e 10.0 53.3 63.4NOLTN – – – – – – – – – –LTN1 10 5 1.0 5.0 176 1447 8.21e 1.8 7.2 9.0LTN2 10 5 1.0 4.8 254 702 2.76 2.5 3.5 6.1LTNWP 67f 6.7 1.0 4.8 254h 702 2.76 17.0 4.7 21.7LTNHWP 300g 30 1.0 4.8 254 702 2.76 76.2 21.1 97.3

a 1025 molecules per flash.b Number of flashes calculated for the 7-day TOGA COARE runs using Eq. (2).c Ratio of number of IC flashes to number of CG flashes.d Total NO production (1028 molecules).e Diagnosed from calculated total flash numbers.f Same as in LTN run ofWang and Prinn(2000), based onPrice et al.(1997) who suggested 6.7·1026 NO

molecules per CG flash.g Same as the LTNH case ofWang and Prinn(2000), based onFranzblau and Popp(1989).h For comparison:Wang and Prinn(2000) calculated a total of about 750 flashes, out of which∼210 were

CG flashes for the first hours of a 30 h 2-D run with 1000 km horizontal domain length.

flash segments are assumed to have a Gaussian distribu-tion and IC segments are assumed to have a bimodal dis-tribution corresponding to a superposition of two Gaussiandistributions. The pressure dependence of the NO produc-tion is taken into account. In the present study, the uppermode of the IC distribution is assumed to be centered atz=z(−15) + 0.8·(zctw−z(−15)), wherez(−15) is the altitude ofthe –15◦C isotherm, thus allowing the altitude of the uppermode to vary depending on the growth stage of the cloud.

CG flashes are placed (horizontally) in the vertical gridcolumn at the location of the maximum updraft velocity,which is consistent withRay et al.(1987), who based ondual Doppler radar and VHF lightning observations foundthat in a multi-cell storm, lightning tended to coincide withthe reflectivity and updraft core. Nevertheless, this assump-tion potentially leads to a small over-estimate of the upwardtransport of lightning NOx. IC flashes are horizontally placedinside the previously identified anvil area in the vertical col-umn whereqtotm abovez(−15) is largest and which is stilllocated within 80 km of the maximum updraft.

For each system, the anvil is identified as follows: First allgrid points in the entire domain with a cloud top heightzct

betweenzctw–1400 m andzctw+1400 m are flagged. Then,one grid point wide connections between flagged regions(“bridges”) and single points (“bumps”) on the edges of theflagged regions are removed. Finally, all the flagged pointsare identified which lie inside the region in which the columnwith the maximum updraft velocity is located, and which areless than 80 km away from this column (see Sect. S1.4 fordetails). This fairly simple method allows us to assign anvils

to the convective systems which have previously been iden-tified by fitting rectangles as described above.

In the 2-D runs lightning-produced NO is released intovertical columns of the same horizontal base area as in the3-D run, i.e., 2 km×2 km. For numerical efficiency, light-ning NO production is calculated every 56 s in the 3-D runand every 60 s in the 2-D runs. This choice to some extentmimics the discreteness of flashes at times when flash ratesare low.

3 Lightning NO sensitivity runs

3.1 Flash rates

Lightning in the TOGA COARE region is relatively infre-quent. Orville et al. (1997), for example, found the annualnumber of CG flashes per unit area in the TOGA COAREregion for 1993 to be at least one order of magnitude lowerthan that observed close to the nearby Papua New Guineanislands. From a number of dedicated observations duringTOGA COARE it is possible to estimate typical CG flashrates per storm and area densities of CG flashes. These es-timates have been used to adjust the parameters in the flashrate parameterization Eq. (2). Large uncertainty exists, how-ever, regarding the ratioZ=NIC/NCG. Furthermore, thenumber of molecules produced per flash is still uncertain.Here, we conduct a number of 2-D sensitivity runs with var-ious assumptions regarding lightning NO production, in ad-dition to a 3-D reference run (see Table1). In the LTN3Dreference run and the LTN1 run,Z is calculated using the



LTNHWP

LTN3D, LTN1, LTN2

LTNWP

Fig. 1. Lightning NO production rates in molecules per flash fromthe literature based on Table 21 ofSchumann and Huntrieser(2007).

parameterization in Eq. (3). In the other runs,Z=2.76 isprescribed for clouds with1h≥5.5 km. The same value ofZ has also been adopted byWang and Prinn(2000) in theircloud resolving model study of a storm during CEPEX. Fur-thermore, in the LTNWP and the LTNHWP run we prescribethe same number of molecules per IC and per CG flash asWang and Prinn(2000) in their LTN and LTNH sensitivityrun, respectively. In our LTN3D, LTN1, and LTN2 run weadopt a lower NO production per CG flash (Fig.1), but agreater ratio of molecules per IC flash to molecules per CGflash based on recent literature (Gallardo and Cooray, 1996;Cooray, 1997; DeCaria et al., 2000; Fehr et al., 2004).

Figure2a shows modeled 30 min flash rates for the LTN3Drun. The area density and maximum number of CG flashesin the LTN3D run are in line with the available observations.The daily area flash density in the TOGA COARE region canroughly be estimated to equal 0.0022 CG flashes per km2 perday from Fig. 10 ofPetersen et al.(1996) on days with highlightning activity during seasons with high lightning activ-ity. The number of CG flashes in our LTN3D run is 1021(Table1) and the average area flash density during the last6 days of the simulation (after the onset of deep convectionin the model) is 0.0028 CG flashes per km2 per day, whichagrees with the observations within∼30%.

In selecting the parameters in the flash rate parameteriza-tion (Eq.2) we attempted to approximatively match observedCG flash area densities as well as flash rates per storm. In se-vere continental thunderstorms rates of 20 and more flashesper minute can be sustained over an hour or more (see e.g.Skamarock et al., 2003; Fehr et al., 2004). In the TOGACOARE region, on the other hand, during more active stormsflash rates of only 1–2 (and in one case 4) discharges (bothIC and CG) per minute were observed by two NASA air-craft during storm overpasses (Orville et al., 1997). Light-ning was only observed during 19 out of 117 storm over-

Fig. 2. (a)Flashes per 30 min in the LTN3D run and smoothed (5point running mean) total flash rates (green line).(b) MaximumNO (solid line) and NOx (dotted line) volume mixing ratios below∼16 km for the LTN3D run. Local ’solar’ time is UTC plus 10.5 h.

passes.Petersen et al.(1996) also noted that the majority ofconvection over the tropical oceans does not produce light-ning. For four cells with similar characteristics which wereall observed on 15 February 1993 they found CG flash ratespeaking at slightly above 7 CG flashes within 20 min (seetheir Fig. 5). Here, the maximum CG flash rates are typi-cally between 10 and 20 flashes per 30 min in all lightningsensitivity runs (see Fig.2a for the LTN3D, Fig.3a for theLTN1 run, and Fig.4a for the other 2-D lightning sensitivityruns). Often only the single most vigorous updraft signifi-cantly contributes to these rates (not shown).

3.2 Lightning NOx

The maximum NO mixing ratios in the LTN3D run (Fig.2b)fall largely within the range of available observations fromthe literature; several airborne observations of deep convec-tive outflow in the Central and West Pacific yielded max-imum NO mixing ratios of slightly below 1 nmol mol−1

(Huntrieser et al., 1998; Kawakami et al., 1997). In the out-flow from continental thunderstorms, on the other hand, max-imum NO mixing ratios of a few nmol mol−1 are not uncom-mon (e.g.Huntrieser et al., 1998). Only occasionally, theoutputs from the LTN3D, LTN1, and LTN2 runs (Figs.2b,3b, and4b) yield higher maxima than those observed overthe Pacific. In the LTNWP and the LTNHWP (Figs.4c andd) run we often find significantly higher maxima than ob-served. Somewhat higher maxima in the model are likely re-lated to aircraft typically sampling deep convective outflowoutside the region of maximum lightning activity. The ex-tremely high NO maxima due to lightning in the LTNHWP



Fig. 3. Same as Fig.2 for the LTN1 run.

run can nevertheless be considered unrealistic.Lower domain-average NOx and NO volume mixing ra-

tios in our LTNWP and LTNHWP runs compared to the vol-ume mixing ratios at the end of the LTN and LTNH runs ofWang and Prinn(2000) (also included in Fig.5) are in partcaused by lower average flash rates in our study and the useof different model setups.Wang and Prinn(2000) used pe-riodic lateral boundary conditions and calculated a total ofabout 750 flashes (estimated from their Fig. 6) out of which∼210 were CG flashes for a 30 h run with a 1000 km 2-Ddomain; i.e. 0.17 CG flashes km−1 day−1 (compared to 0.07CG flashes km−1 day−1 in the LTN2 run).

The scaling factorα in Table 1 is smaller for the 3-D runthan for the 2-D run because of generally higher vertical ve-locities in 3-D simulations compared to 2-D simulations (Re-delsperger et al., 2000; Phillips and Donner, 2006). In spiteof the smallerα in the LTN3D run, the number of flashesand the number of NO molecules produced are greater in theLTN3D run than in the LTN1 run. Because of the differ-ent geometry, the domain-averaged mixing ratios in Fig.5are, nevertheless, higher in the LTN1 run; i.e. by assuminglightning-produced NO to be released into vertical columnsof the same horizontal base area (2 km×2 km) in the 2-D runsas in the 3-D run, it is implicitly assumed that the horizontal“area” of the 2-D domain equals 2 km×500 km. For a givennumber of NO molecules produced by lightning, this leads togenerally greater domain-average NOx volume mixing ratiosdue to lightning in the 2-D simulations. Assuming a lowerNO production per flash in the 2-D runs in order to compen-sate for this geometry effect, on the other hand, would alsolead to unrealistic consequences for chemistry in the vicinityof the lightning NO source.

Fig. 4. (a) Same as Fig.2a for the LTN2, the LTNWP, and theLTNHWP run. (b-d) Same as Fig.2b for the LTN2, the LTNWP,and the LTNHWP run, respectively.

Using a vertical distribution of NO molecules which isconstant except for being scaled by pressure for IC and CGflashes instead of theDeCaria et al.(2000) parameteriza-tion results in smaller maxima for the same number of NOmolecules produced per flash (not shown), because of the twomaxima in the vertical distribution byDeCaria et al.(2000)(compare also “lno” term in Fig.8b). The influence of light-ning on the average NOx mixing ratio profiles is small in our3-D run, but considerable in the 2-D runs (Fig.5). An exam-ple of the spatial distribution of NOx from the LTN3D runis shown in Fig.6. An example from a 2-D run is shownin Fig. 7. Note, that this NOx plume is almost completelyadvected out of the domain in∼2.5–3 hours.

3.3 NOx budget

In order to assess the influences of deep convection on chem-istry in the model simulations we performed budget calcula-tions. For each tendency on the right hand side of Eq. (1) we



define height dependent contributions by:

bproc(z, t1, t2) =1

XY

∫ X

0

∫ Y

0

∫ t2

t1

∂tCi,g|proc dx dy dt (4)

whereX and Y is the horizontal extent of the domain inWest-East and South-North direction, respectively. Horizon-tal advection∂tCi,g|hadv=−∇ ·

(vhCi,g

)and vertical advec-

tion ∂tCi,g|vadv=−∂z

(wCi,g

)(wherevh=(u, v) is the hori-

zontal wind vector) are considered separately. The budgetterms have been divided by time-averaged air density in or-der to express them in mixing ratio units. In addition to the“height resolved” terms in Eq. (4), the corresponding heightintegrated terms are analyzed for layers betweenz1 andz2:

Bproc=

∫ z2

z1

bprocdz (5)

Figure8a shows NOx mixing ratio profiles for the LTN3Drun, and Fig.8b shows time integrated domain-averaged ten-dencies. Most of the NOx produced below 10 km is trans-ported to the upper troposphere from where it is horizontallyadvected out of the domain (compare also Fig.6 and Fig.7).The lightning production term has two maxima reflecting thevertical distribution functions fromDeCaria et al.(2000),in which IC as well as CG flashes contribute to the mid-tropospheric maximum while the upper tropospheric maxi-mum is caused by IC flashes. Very little NOx reaches thelowest kilometer of the atmosphere, a region where detrain-ment from deep cloud downdrafts can be expected. This re-sult is consistent with a result ofPickering et al.(1998) froma CRM study of a single storm during TOGA COARE, whofound only about 7.5% of the lightning-produced N mass inthe lowest km because of the very weak downdrafts in theirsimulated storm. In mid-latitude storms they found consider-ably larger fractions. In the CSRMC, the downward transportcould be underestimated because lightning NO is assumed tobe produced in columns with maximum updraft velocities.

Net photochemical loss of NOx within the domain playsa relatively much more important role in our 2-D lightningsensitivity runs than in our 3-D run (Fig.9). Higher NOxvolume mixing ratios in the 2-D runs lead to increases ofOH concentrations (Sect.5) and a reduced chemical lifetimeof NOx. The fraction of NOx photochemically lost withinthe domain varies between 20 and 24% in the 2-D runs, butis only 1% in the LTN3D run. In the 3-D run the loss oflightning-produced NOx in the lower troposphere is in factover-compensated by production from PAN. The net chem-ical production of NOx is 2.0×1013 molecules cm−2 day−1

in the lower troposphere (LT, 0–5 km) during the last 6 daysof the simulation, when deep convection is active. Byfar the largest chemical source of NOx in the LT is PAN,while HNO3 is the largest sink. PAN is horizontally ad-vected into the domain above∼3 km. More details on theNOx, PAN, and HNO3 budgets can be found in the supple-mentary materialhttp://www.atmos-chem-phys.net/8/2741/2008/acp-8-2741-2008-supplement.pdf.

Fig. 5. Domain- and time-average (black lines) and median (redlines) volume mixing ratios of(a) NOx and(b) NO for solar zenithangles (SZA) lower than 70◦. Blue lines in (a): domain-averagedNOx volume mixing ratio from Fig. 11 ofWang and Prinn(2000)at the end of their LTN and LTNH runs.

Comparable 2-D sensitivity runs with a pressure scaledbut otherwise constant vertical distribution of NO moleculesfor IC and CG flashes yielded different fractions of NOxlost within the domain (13–37%) and greater sensitivitiesto various assumptions regarding lightning NOx production(Fig. 8.8 ofSalzmann, 2005). In a non-HOx-limited regime,the relatively smaller photochemical loss in the LTN3D runis in agreement with the lower average NOx concentrationscompared to the 2-D run. Furthermore, the shorter residencetime in the 3-D domain probably also plays a role, but this isdifficult to quantify.

4 Ozone

The sensitivity of the domain-averaged O3 mixing ratios(Fig. 10a) to in situ lightning NO production is small anda moderate increase of ozone volume mixing ratios withincreasing lightning NO production is simulated only inthe mid-troposphere. The chemistry tendencies (shown inFig. 10b) in percent per day in the NOLTN and the LTN3Drun are similar to the ones calculated byCrawford et al.(1997) for their “low NOx regime” (Table S6). They reflectthe influence of lightning on ozone within the domain nottaking into account ozone production further downwind. Inthe lower troposphere, the simulated O3 mixing ratios arealmost independent of the assumptions regarding lightningNOx production, even in the sensitivity runs with very high





Fig. 6. Spatial distribution of NOx in the LTN3D run for 24 December 1992, 10:30 UTC: 50 pmol mol−1 (dark gray, transparent) and100 pmol mol−1 (light gray) isosurfaces, surface to 16 km column density contours (base area), and average volume mixing ratio in x (East-West) and y (South-North) direction (walls).

NOx production by CG flashes. While here the sensitivity oflocal O3 mixing ratios to in situ lightning NOx production issmall, Crawford et al.(1997) found considerably increasedO3 mixing ratios in air masses sampled during PEM-West Bwhich had presumably been influenced by lightning overland upstream. These air masses had been more chemicallyaged and better mixed.

4.1 Ozone budget

The domain-averaged O3 mixing ratios vary significantlyduring the LTN3D run (Fig.11a) and are very similar to theboundary values (Fig.11b), indicating that horizontal advec-tion from the lateral boundaries rapidly influences and ulti-mately determines the simulated domain averages. Althoughvertical advection (Fig.11e) tends to balance horizontal ad-vection (Figs.11d and12b), the total tendency (Fig.11c)mainly reflects horizontal advection from the boundaries.The time integrated total O3 tendency (equaling1O3 inFig. 12), on the other hand, is smaller than the time inte-grated contributions from either horizontal or vertical advec-tion, i.e., over longer time scales an approximate equilib-rium (1O3≈0) tends to establish. Furthermore, the contri-bution from photochemistry plays an important role in the7-day integrated domain budget (Fig.12). The most pro-nounced net loss takes place at∼5 km where relatively O3-

rich air is advected into the domain and where photolysisrates are occasionally enhanced through sunlight reflected bymid-level cloud tops. While vertical advection of O3-poorair on average tends to decrease O3 volume mixing ratiosabove∼14 km, vertical advection of ozone from the mid-troposphere (between∼6–10 km) enhances O3 mixing ratiosbetween∼10 and 14 km. This is consistent with our previousstudy of idealized tracer transport (Salzmann et al., 2004) inwhich we found considerable mid-tropospheric entrainmentin a 3-D TOGA COARE simulation.

Horizontal advection plays a somewhat weaker role in de-termining the domain-average O3 mixing ratios in the 2-Druns (Fig.13) because of the larger 2-D domain and becausemeridional advection is not taken into account in the 2-Druns. Furthermore, in the LTN1 run the advection of O3-poor air from below tends to decrease O3 mixing ratios in theupper troposphere between∼12 and 16 km (Fig.14). Theless obvious influence of mid-level entrainment in the 2-Drun can be attributed to the weaker dependence on the lateralboundary values, i.e., horizontal layers of high O3 mixingratios in the mid-troposphere are more efficiently diluted bymixing in the 2-D runs. This is because in the 3-D run up-drafts only cover a small part of the domain and only a partof the area downwind of the updrafts is directly affected. In2-D runs, on the other hand, a single updraft influences thewhole downwind part of the domain, and a significant frac-



Fig. 7. NOxvolume mixing ratios and theqtotm=0.01g kg−1 con-tour (whereqtotm=qcloud water+ qrain + qice + qgraupel+ qsnowis the sum of all hydrometeor mass mixing ratios) during the de-velopment of a mesoscale convective system from 24 December1992, 10:00 UTC to 24 December 1992, 12:30 UTC from the LTN1run. The time increment between the plots in each of the panels is30 min.

Fig. 8. (a)Time- and domain-averaged NOx volume mixing ratio,averaged NOx volume mixing ratio± standard deviation, initial,and final NOx volume mixing ratio profiles for the LTN3D run.(b)Budget terms (from Eq. (4), divided by the average air density) andthe difference between the initial and the final profile (red line).

Fig. 9. Ratios of time integrated tropospheric column NOx tenden-cies for the lightning sensitivity runs (calculated from Table S4).

Fig. 10. (a)Domain- and time-averaged (black) and median (red)O3 volume mixing ratio for SZA<70◦. (b) Domain- and time-averaged net O3 chemistry tendency for SZA<70◦.

tion of the updrafts subsequently transports O3-poor air tothe upper troposphere. These updrafts do not pass throughthe mid-tropospheric high ozone layer, while in a 3-D sim-ulation practically all updrafts can be expected to encountersuch a layer at times when O3-rich air is advected into the do-main. Due to the greater fraction of updrafts in the 2-D runswhich do not entrain O3-rich air from the mid-troposphere,deep convective transport is on the whole less strongly af-fected by these layers in 2-D runs than in a 3-D run (com-pare Figs. 8c and 12b ofSalzmann et al., 2004). In spite ofthese differences in transport, the lower tropospheric chem-istry tendency is very similar in the LTN1 and the LTN3Drun (Fig.10b).



Fig. 11. (a) Time series of domain-averaged O3 volume mixingratios for the LTN3D run;(b) boundary values;(c)–(f) budget termsas defined by Eq. (4).

4.2 Undiluted transport

Global and also mesoscale models often use bulk instead ofensemble mass flux parameterizations in order to calculatedeep convective tracer transport, in which mid-level entrain-ment is incompatible with undiluted transport to the uppertroposphere (Lawrence and Rasch, 2005). In Fig. 15, simu-lated upper tropospheric O3 volume mixing ratio minima arecompared to minima in the lowest model layer. The UT min-imum in the LTN3D run is often much closer to the surfaceminimum than in the 2-D run, indicating that some undilutedair has reached the upper troposphere in this run in spite ofsignificant mid-level entrainment (Sect.4.1).

During the CEPEX cruise, extremely low (<5 nmol/mol)upper tropospheric O3 volume mixing ratios were observed(e.g.Kley et al., 1996, 1997). A few model studies (Wanget al., 1995; Lawrence et al., 1999a) have reproduced theseminima qualitatively, but not quantitatively. An exceptionhas been the cloud resolving model study byWang and Prinn(2000), who suggested that lightning related O3 loss can helpexplain layers of very low O3 in the upper tropical tropo-sphere. Figure15shows pronounced decreases of O3 mixingratios indicating the loss of O3 in the reaction with NO onlyfor the LTNHWP run and to a much lesser extent also forthe LTNWP run. Since the number of molecules per flash inthe LTNHWP run is several times as large as the next largest

Fig. 12. As Fig.8 for ozone.

Fig. 13. As Fig.11 for the LTN1 run.

estimate from the literature (Fig.1), it is extremely unlikelythat the occasionally observed layers of very low O3 concen-tration are caused by lightning NO production.

In the present study a reproduction of the extremely lowupper tropospheric O3 concentrations observed during theCEPEX cruise was not to be anticipated since near-zerolower tropospheric O3 volume mixing ratios can mainly beexpected during the easterly phase of the intra-seasonal os-cillation (see discussion in Sect.6). Furthermore, the lifetimeof O3 in the MBL is of the order of few days and the lat-eral boundary conditions do not yield very low O3 concentra-tions in either the upper or the lower troposphere. In order to



Fig. 14. As Fig.12 for the LTN1 run.

reproduce the very low ozone concentrations observed dur-ing CEPEX, one would most likely need a larger domain.Furthermore, as has been suggested inKley et al. (1996);Lawrence et al.(1999a), it might be necessary to include ad-ditional reactions involving halogens (e.g.Davis et al., 1996;Sander and Crutzen, 1996, compare discussion in Sect.6).

4.3 Small scale downward transport

Small scale downward transport of air with stratospheric ori-gin in association with deep convection has been simulatedin a number of two-dimensional cloud resolving model stud-ies (Scala et al., 1990; Wang et al., 1995; Stenchikov et al.,1996; Wang and Prinn, 2000) in agreement with observa-tions by Poulida et al.(1996) and Dye et al.(2000), whofound air with stratospheric characteristics in the upper tro-posphere in the vicinity of mid-latitude storms. A very pro-nounced example from our LTN1 run is shown in Fig.16where O3-rich air is advected downward in the rear inflowof a large mesoscale convective system. In the LTN3D run,such events are less easily identified, in part possibly due tothe lower vertical resolution.Newell et al.(1996, 1999) inter-preted thin horizontal layers in the troposphere as frozen sig-natures of the transfer of boundary layer air into the free tro-posphere or of stratospheric air into the troposphere.Newellet al. (1996) identified over 500 layers with a mean thick-ness of about 400 m from 105 vertical profiles sampled dur-ing PEM-West A over the West Pacific. They found theselayers to extend over distances of 100–200 km or more. Us-ing a different method to identify horizontal layers,Newellet al. (1999) found the average thickness of horizontal lay-ers during PEM-West A to be about 800 m. The verticalresolution in this study is not high enough to adequatelyresolve the thinner ones among these layers. Instead, wefocus on a comparison withWang and Prinn(2000), whofound that downward transport of O3-rich air increased up-per tropospheric O3 mixing ratios, outweighing the effectof upward transport of O3-poor air. In the LTN1 run, onthe other hand, vertical advection tends to decrease upper

(a)

(b)

Fig. 15. Time series of minimum O3 mixing ratios in the lowestmodel layer (dashed lines) and in the UT (defined from about 10 kmto about 16 km, dotted lines) for(a) the LTN3D run and(b) the 2-Dlightning sensitivity runs.

tropospheric O3 mixing ratios (Fig.14). On the whole, asmall net amount of O3 is transported from below 16 km(troposphere) to above 16 km (stratosphere) in our simula-tions: 4×1015 molecules cm−2 day−1 in the LTN3D run, and6×1015 molecules cm−2 day−1 in the LTN1 run. This smallupward transport is influenced by the method used to spec-ify the large scale forcing (see Sect. S5.3 for details) and isconsistent with parts of the tropics being a region of averagelower stratospheric ascent and troposphere to stratospheretransport (Holton et al., 1995; Corti et al., 2005).

The reason whyWang and Prinn(2000) found downwardtransport to increase upper tropospheric O3 mixing ratiosin spite of the upward transport of O3-poor air might bestbe explained by their initial profile, which varied by only∼3–4 nmol mol−1 between the surface and the upper tropo-sphere, so that the upward transport of O3-poor air playeda relatively smaller role. In order to better quantify the in-fluences of small scale downward transport on upper tro-pospheric O3 concentrations in the tropics, additional de-tailed observations of O3 and anthropogenic pollutants in



Fig. 16. O3 volume mixing ratios (filled) from the LTN1 run for24 December 1992, 9:30 UTC to 13:30 UTC and total hydrometeormixing ratioqtotm=0.01g kg−1 contour.

the vicinity of deep convection are necessary that allow usto discriminate between vertical transport from below andimport from the stratosphere, in combination with dedicated3-D model studies, possibly using an idealized O3 tracer andvery high vertical resolution.

5 HOx

Together with NOx, hydrogen oxides play a crucial role forthe troposphere’s oxidation capacity and ozone chemistry.Figure17 shows increasing OH concentrations with increas-ing lightning NO production in the 2-D lightning sensitivityruns and slightly decreasing HOx concentrations in all runsbut the LTNHWP run, in which the decrease is somewhatlarger (of the order of a few pmol mol−1 in the upper tropo-sphere). An increase of OH and a small decrease of HO2 iscaused by the NO oxidation reaction

NO + HO2 → NO2 + OH

Fig. 17. (a)Median (red) and average (black lines) OH concentra-tion around local noon and diurnal averages (blue).(b) Same forHOx volume mixing ratios.

which shifts the partitioning between OH and HO2 to-wards higher OH/HO2 ratios while leaving the HOx con-centration constant. At very high NO concentrations in thevicinity of the flashes HOx concentrations decrease due tothe termination reaction of OH with NO2 in which HNO3 isformed. While we find increasing OH concentrations withincreasing lightning NO production,Wang and Prinn(2000)found a very strong decrease of OH in one of their lightningsensitivity runs and a complete depletion of OH throughoutthe troposphere in the other run, which is consistent withsignificantly higher NOx concentrations in their simulations(Fig. 5). The elevated HOx concentrations above the bound-ary layer in Fig.17 are caused by larger scale horizontaltransport of O3-rich air (compare Sect.4 and Sect.6). Ele-vated OH concentrations above the boundary layer have pre-viously been observed above the eastern tropical Pacific un-der conditions of high NOx and CO concentrations at thesealtitudes (Davis et al., 2001).

6 Role of the intra-seasonal oscillation (ISO) for ozoneprofiles

Kley et al.(1996, 1997) found extremely low O3 mixing ra-tios during CEPEX either in the lower or the upper tropo-sphere (depending on wind direction), but not in the upperand the lower troposphere at the same time. Figure18 sug-gests that the surface concentrations of O3, CO, and NOxfrom MATCH-MPIC are strongly influenced by the phaseof the intra-seasonal (Madden-Julian) oscillation (ISO, e.g.Madden and Julian, 1994) as reflected by the 10 m zonalwind. During the westerly phase (WP), relatively O3, CO,and NOx-rich air is found in the lowest model layer. Whilethe zonal wind velocity in the MBL (below∼1000 m) ispositively correlated with minimum O3 volume mixing ra-tios in the MBL just south of the Equator between 150◦Eand 180◦E (Fig. 19a), upper tropospheric minimum O3 vol-ume mixing ratios are negatively correlated with the zonalwind velocity in the MBL (Fig.19b); i.e., during the WP,



Fig. 18. Time series of(a) 10 m observed zonal wind,(b)10 m zonal wind from the interpolated NCEP re-analysis data (seeSect. S8),(c) O3 volume mixing ratio in the lowest model layer(surface layer) calculated by MATCH-MPIC,(d) surface layer CO,(e)surface layer NOx at the site of the TOGA COARE IFA.

O3 volume mixing ratios in the MBL tend to be higher andupper tropospheric O3 volume mixing ratios lower than dur-ing the easterly phase. The zonal wind component lags O3in the MBL (Fig. 19a, lower panel), but taking into accounta lag does not improve the correlation significantly at mostlongitudes. The increase in lower tropospheric O3 volumemixing ratios during the WP is most likely related to cross-equatorial flow over the Indonesian Archipelago advectingpolluted air masses from the Northern Hemisphere into theTOGA COARE/CEPEX region (see schematic in Fig.20).This cross-equatorial flow is characteristic of the westerlyphase of the ISO (e.g.Yanai et al., 2000; McBride et al.,1995). While surface winds are westerly during the WP, up-per tropospheric winds tend to be easterly, advecting O3-poorair from convective regions of the central Pacific into the re-gion. During the easterly phase of the ISO the flow is roughlyopposite. The dependence of O3 concentrations on the phaseof the ISO and regional scale horizontal advection can help toexplain the lack of coincidence between upper troposphericand lower tropospheric O3 minima in the profiles observedduring CEPEX. Furthermore, one could speculate that thelow (high) NOx regime (in the upper troposphere) identifiedby Crawford et al.(1997) might coincide with the westerly(easterly) phase of the ISO. The seven-day episode from 19–26 December 1992 is characterized by the onset of a strongWP.

Fig. 19. Upper panels: Correlation of vertically averaged zonalwind velocity below∼1000 m to(a) minimum O3 volume mixingratios in the MBL (below∼1000 m) and(b) minimum O3 volumemixing ratios in the upper troposphere (∼10–15 km), at 2.86◦S for31 October 1992 to 31 March 1993 from MATCH-MPIC for lag 0 h(diamonds) and for a lag between –96 and 96 h (lower panels) at thetime of the maximum absolute correlation (crosses).

MATCH-MPIC reproduces the ozone minima observedduring CEPEX qualitatively but not quantitatively (Lawrenceet al., 1999a); a possible candidate for a missing chemicalozone sink in the model being halogen-catalyzed ozone de-struction (e.g.Davis et al., 1996; Sander and Crutzen, 1996).

The domain of the CSRMC is too small to simulate theregional scale transport associated with the ISO explicitly;instead the effect of regional scale horizontal advection istaken into account by specifying boundary conditions fortrace gases. In the future, growing computer power will allowus to increase domain sizes and a multiply-nested cloud re-solving model setup including chemistry for TOGA COAREwill become feasible. A nested setup without chemistry iscurrently being evaluated for TOGA COARE.

7 Summary and conclusions

A cloud system resolving model including a simple light-ning NO parameterization has been developed which allowsus to study the influences of multiple storms and mesoscaleconvective systems on tropospheric chemistry in the TOGACOARE/CEPEX region over a period of several days. In situlightning NO production is simulated to have a minor impact



UT

LT

Fig. 20. Schematic: During the westerly phase of the ISO lower tropospheric cross-equatorial flow transports polluted air to the TOGACOARE/CEPEX region. The contours at the base are 850 hPa ozone volume mixing ratios from MATCH-MPIC and interpolated windvectors are from the NCEP re-analysis dataset for 24 December 1992, 00:00 UTC.

on domain-averaged O3, and the lightning NOx plumes arehorizontally advected out of the domain. The effect of light-ning NO on domain-averaged NOx strongly depends on thedimensionality of the model. In the 3-D reference (LTN3D)run lightning has a weak influence on the domain-averageNOx profile, but in the 2-D sensitivity runs lightning NOcauses a large increase of domain-averaged NOx because ofthe smaller implicitly assumed “volume” of the 2-D domain.Significantly lower NOx volume mixing ratios in the LTN3Drun, in spite of a much higher number of flashes, have animportant implication for the loss of NOx in the vicinity ofthe storms: While between 20 and 24% of the lightning-produced NOx is lost due to chemistry inside the domain inthe 2-D runs, this fraction is negligible in the LTN3D runand the loss is more than compensated by NOx productionfrom PAN. This suggests that the results of 2-D sensitivityruns are to be interpreted with caution and that 3-D runs areneeded in order to realistically assess the influences of in situlightning on local chemistry. Because multi-day 3-D runs arestill computationally expensive (to our knowledge this is thefirst multi-day cloud system resolving study including chem-istry), 2-D sensitivity runs can nevertheless be useful, espe-cially for comparing with 2-D sensitivity runs from a previ-ous study. Domain-averaged OH is found to increase withincreasing lightning NO production and HOx is fairly insen-sitive except for a notable decrease in the run with extremelyhigh NO production per flash.

Mid-level entrainment plays an important role for O3transport from the mid-troposphere to the UT between∼10and 14 km, but transport of O3-poor air from the MBL tendsto decrease O3 volume mixing ratios between 14 and 15.5 kmin the LTN3D run. In spite of the importance of mid-levelentrainment, undiluted O3-poor air from the MBL reachesthe upper troposphere more frequently in this run than in

the 2-D runs, indicating the presence of undiluted convec-tive cores (e.g.Heymsfield et al., 1978). Unlike in a bulkupdraft cumulus transport parameterization or single columnmodel, mid-level entrainment does not preclude the trans-port of undiluted air in a cloud system model since the dy-namics of multiple storms and mesoscale convective systemsis modeled. The finding of undiluted transport despite mid-level entrainment strongly suggests that plume ensemble pa-rameterizations (e.g.Arakawa and Schubert, 1974; Gidel,1983; Donner, 1993; Lawrence and Rasch, 2005) should beused instead of bulk mass flux parameterizations in globalchemistry-transport models.

The causes of extremely low O3 mixing ratios in the up-per troposphere were examined through a series of sensitivitystudies. We found that low O3 mixing ratios can be a result oflightning-produced NO when high NO production rates andflash rates are used in a 2-D run, similar toWang and Prinn(2000). However, a recent review of NO production by light-ning (Schumann and Huntrieser, 2007) indicates that the NOproduction rate used in this sensitivity simulation is 10–100times greater than most values reported in the literature. Wetherefore think this source is unrealistic.

We find vertical transport of ozone poor air to play a largerrole in causing low upper tropospheric mixing ratios thanWang and Prinn(2000), who specified an initial O3 profilewith a relatively small vertical gradient between the bound-ary layer and the upper troposphere.

A reproduction of the extremely low O3 concentrations ob-served during CEPEX (which started in the TOGA COAREregion) was not to be anticipated in the present study sinceduring the westerly phase of the intra-seasonal oscillation(ISO), relatively O3-rich air is transported to the regionby low level westerlies. Correlations between zonal windand ozone volume mixing ratios from MATCH-MPIC for



31 October 1992 to 31 March 1993 suggest that verticalozone profiles in the TOGA COARE/CEPEX region arestrongly influenced by the ISO. This influence can help toexplain the lack of coincidence between upper and lower tro-pospheric O3 minima observed during CEPEX.

In the LTN1 run, small scale downward transport of ozonefrom the tropopause layer is simulated to occur in thin fila-ments extending far into the troposphere in association withrear inflow into mesoscale convective systems. On the whole,however, vertical advection tends to decrease upper tropo-spheric ozone mixing ratios in this run due to the upwardtransport of O3-poor air from the MBL, and a small aver-age net upward transport of O3 from below∼16 km in as-sociation with average large scale ascent is simulated in allruns. In order to better quantify the influences of small scaledownward transport, additional observations in the vicinityof deep convection are necessary as well as dedicated 3-Dmodel studies with very high vertical resolution. Since cloudsystem resolving model simulations require meteorologicalinput from comprehensive observation campaigns such asTOGA COARE nested cloud resolving model simulationsare likely to be a good alternative in many regions. Further-more, as discussed inSalzmann et al.(2004), the applicationof observation derived large scale forcings for trace gaseswould, unlike for water vapor and potential temperature, beproblematic.

Acknowledgements.We are very thankful to Larry Horowitz forproviding pre-processed NCEP meteorological data for MATCH-MPIC input and to Rolf von Kuhlmann for useful discussions andadvice on the chemistry mechanism and the implementation ofthe ice uptake. Furthermore, we appreciate the comments by tworeferees which have led to many significant improvements of ourinitial manuscript. This research was supported by funding fromthe German Research Foundation (DFG) Collaborative ResearchCentre 641 (SFB 641) “The Tropospheric Ice Phase (TROPEIS)”and from the German Ministry of Education and Research (BMBF),project 07-ATC-02. The model output files in netCDF format(∼180Gb in total) can be obtained upon request from the authors atthe MPI-C.

Edited by: Ronald Cohen

References

Arakawa, A. and Schubert, W. H.: Interaction of a cumulus cloudensemble with the large-scale environment, Part I, J. Atmos. Sci.,31, 674–701, 1974.

Baker, M. B., Christian, H. J., and Latham, J.: A computationalstudy of the relationships linking lightning frequency and otherparameters, Q. J. R. Meteorol. Soc., 121, 1525–1548, 1995.

Christian, H. J., Blakeslee, R. J., Boccippio, D. J., Boeck, W. L.,Buechler, D. E., Driscoll, K. T., Goodman, S. J., Hall, J. M.,Koshak, W. J., Mach, D. M., and Stewart, M. F.: Global fre-quency and distribution of lightning as observed from space by

the Optical Transient Detector, J. Geophys. Res., 108, 4005,doi:10.1029/2002JD002347, 2003.

Ciesielski, P. E., Johnson, R. H., Haertel, P. T., and Wang, J.: Cor-rected TOGA COARE sounding humidity data: Impact on diag-nosed properties of convection and climate over the warm pool,J. Climate, 16, 2370–2384, 2003.

Cooray, V.: Energy dissipation in lightning flashes, J. Geophys.Res., 102, 21 401–21 410, 1997.

Corti, T., Luo, B. P., Peter, T., Vomel, H., and Fu, Q.: Mean ra-diative energy balance and vertical mass fluxes in the equatorialupper troposphere and lower stratosphere, Geophys. Res. Lett.,32, L06802, doi:10.1029/2004GL021889, 2005.

Crawford, J. H., Davis, D. D., Chen, G., Bradshaw, J., Sandholm,S., Kondo, Y., Merill, J., Liu, S., Browell, E., Gregory, G.,Anderson, B., Sachse, G., Barrick, J., Blake, D., Talbot, R.,and Pueschel, R.: Implications of large shifts in troposphericNOx levels in the remote tropical Pacific, J. Geophys. Res., 102,28 447–28 468, 1997.

Davis, D., Crawford, J., Liu, S., McKeen, S., Bandy, A., Thornton,D., Rowland, F., and Blake, D.: Potential impact of iodine on tro-pospheric levels of ozone and other critical oxidants, J. Geophys.Res., 101, 2135–2147, 1996.

Davis, D., Grodzinsky, G., Chen, G., Crawford, J., Eisele, F.,Mauldin, L., Tanner, D., Cantrell, C., Burne, W., Tan, D.,Faloona, I., Ridley, B., Montzka, D., Walega, J., Grahek, F.,Sandholm, S., Sachse, G., Vay, S., Anderson, B., Avery, M.,Heikes, B., Snow, J., O’Sullivan, D., Shetter, R., Lefer, B.,Blake, D., Blake, N., Carroll, M., and Wang, Y.: Marine lati-tude/altitude OH distributions: Comparison of Pacific Ocean ob-servations with models, J. Geophys. Res., 106, 32 691–32 708,2001.

DeCaria, A. J., Pickering, K. E., Stenchikov, G. L., Scala, J. R.,Stith, J. L., Dye, J. E., Ridley, B. A., and Laroche, P.: Acloud-scale model study of lightning-generated NOx in an indi-vidual thunderstorm during STERAO-A, J. Geophys. Res., 105,11 601–11 616, 2000.

DeCaria, A. J., Pickering, K. E., Stenchikov, G. L., and Ott,L. E.: Lightning-generated NOx and its impact on tropo-spheric ozone production: A three-dimensional modeling studyof a Stratosphere-Troposphere Experiment: Radiation, Aerosolsand Ozone (STERAO-A) thunderstorm, J. Geophys. Res., 110,D14303, doi:10.1029/2004JD005556, 2005.

Dickerson, R. R., Huffman, G. J., Luke, W. T., Nunnermacker, L. J.,Pickering, K. E., Leslie, A. C. D., Lindsey, C. G., Slinn, W. G. N.,Kelly, T. J., Daum, P. H., Delaney, A. C., Greenberg, J. P., Zim-merman, P. R., Boatman, J. F., Ray, J. D., and Stedman, D. H.:Thunderstorms: An important mechanism in the transport of airpollutants, Science, 235, 460–465, 1987.

Donner, L. J.: A cumulus parameterization including mass fluxes,vertical momentum dynamics, and mesoscale effects, J. Atmos.Sci., 50, 889–906, 1993.

Dye, J. E., Ridley, B. A., Skamarock, W., Barth, M., Venticinque,M., Defer, E., Blanchet, P., Thery, C., Laroche, P., Baumann,K., Hubler, G., Parrish, D. D., Ryerson, T., Trainer, M., Frost,G., Holloway, J. S., Matejka, T., Bartels, D., Fehsenfeld, F. C.,Tuck, A., Rutledge, S. A., Lang, T., Stith, J., and Zerr, R.: Anoverview of the Stratospheric-Tropospheric Experiment: Radi-ation, Aerosols, and Ozone (STERAO)-deep convection experi-ment with results for the July 10, 1996 storm, J. Geophys. Res.,



105, 10 023–10 045, 2000.Fehr, T., Holler, H., and Huntrieser, H.: Model study on production

of lightning-produced NOx in a EULINOX supercell storm, J.Geophys. Res., 109, D09102, doi:10.1029/2003JD003935, 2004.

Franzblau, E. and Popp, C. J.: Nitrogen oxides produced from light-ning, J. Geophys. Res., 94, 11 089–11 104, 1989.

Gallardo, L. and Cooray, V.: Could cloud-to-cloud discharges beas effective as cloud-to-ground discharges in producing NOx?,Tellus, 48B, 641–651, 1996.

Ganzeveld, L. and Lelieveld, J.: Dry deposition parameterizationin a chemistry general circulation model and its influence onthe distribution of reactive trace gases, J. Geophys. Res., 100,20 999–21 012, 1995.

Gidel, L. T.: Cumulus cloud transport of transient tracers, J. Geo-phys. Res., 88, 6587–6599, 1983.

Heymsfield, A. J., Johnson, P. N., and Dye, J. E.: Observations ofmoist adiabatic ascent in northeast Colorado cumulus congestusclouds, J. Atmos. Sci., 35, 1689–1703, 1978.

Holton, J. R., Haynes, P. H., McIntyre, M. E., Douglass, A. R.,Rood, R. B., and Pfister, L.: Stratosphere-troposphere exchange,Rev. Geoph., 33, 403–439, 1995.

Huntrieser, H., Schlager, H., Feigl, C., and Holler, H.: Transportand production of NOx in electrified thunderstorms: Survey ofprevious studies and new observations at midlatitudes, J. Geo-phys. Res., 103, 28 247–28 264, 1998.

Kawakami, S., Kondo, Y., Koike, M., Nakajima, H., Gregory, G. L.,Sachse, G. W., Newell, R. E., Browell, E. W., Blake, D. R., Ro-driguez, J. M., and Merrill, J. T.: Impact of lightning and convec-tion on reactive nitrogen in the tropical free troposphere, Geo-phys. Res. Lett., 102, 28 367–28 384, 1997.

Kley, D., Crutzen, P. J., Smith, H. G. J., Vomel, H., Oltmans, S. J.,Grassl, H., and Ramanathan, V.: Observations of near-zero ozoneconcentrations over the convective Pacific: Effects on air chem-istry, Science, 274, 230–233, 1996.

Kley, D., Smit, H. G. J., Vomel, H., Grassl, H., Ramanathan, V.,Crutzen, P. J., Williams, S., Meywerk, J., and Oltmans, S. J.:Tropospheric water-vapour and ozone cross-sections in a zonalplane over the central equatorial Pacific Ocean, Q. J. R. Meteorol.Soc., 123, 2009–2040, 1997.

Labrador, L. J., von Kuhlmann, R., and Lawrence, M. G.: Strongsensitivity of the global mean OH concentration and the tropo-spheric oxidizing efficiency to the source of NOx from lightning,Geophys. Res. Lett., 31, L06102, doi:10.1029/2003GL019229,2004.

Landgraf, J. and Crutzen, P. J.: An efficient method for online cal-culations of photolysis and heating rates, J. Atmos. Sci., 55, 863–878, 1998.

Lawrence, M. G. and Rasch, P. J.: Tracer transport in deep con-vective updrafts: Plume ensemble versus bulk formulations, J.Atmos. Sci., 62, 2880–2894, 2005.

Lawrence, M. G., Crutzen, P. J., and Rasch, P. J.: Analysis of theCEPEX ozone data using a 3D chemistry-meteorology model, Q.J. R. Meteorol. Soc., 125, 2987–3009, 1999a.

Lawrence, M. G., Crutzen, P. J., Rasch, P. J., Eaton, B. E., andMahowald, N. M.: A model for studies of tropospheric photo-chemistry: Description, global distributions, and evaluation, J.Geophys. Res., 104, 26 245–26 277, 1999b.

Lawrence, M. G., von Kuhlmann, R., Salzmann, M., andRasch, P. J.: The balance of effects of deep convective mix-

ing on tropospheric ozone, Geophys. Res. Lett., 30, 1940,doi:10.1029/2003GL017644, 2003.

Lelieveld, J. and Crutzen, P. J.: Role of deep cloud convection inthe ozone budget of the troposphere, Science, 264, 1759–1761,1994.

Madden, R. A. and Julian, P. R.: Observations of the 40–50 daytropical oscillation – A review, Mon. Weather Rev., 122, 814–837, 1994.

Mahowald, N. M., Rasch, P. J., and Prinn, R. G.: Cumulus parame-terizations in chemical transport models, J. Geophys. Res., 100,26 173–26 189, 1995.

McBride, J. L., , Puri, K., Davidson, N. E., and Tyrell, G. C.: Theflow during TOGA COARE as diagnosed by the BMRC tropicalanalysis and prediction system, Mon. Weather Rev., 123, 717–736, 1995.

Newell, R. E., Wu, Z.-X., Zhu, Y., Hu, W., Browell, E. V., Gregory,G. L., Sachse, G. W., Collins Jr., J. E., Kelly, K. K., and Liu,S. C.: Vertical fine-scale atmospheric structure measured fromNASA DC-8 during PEM-West A, J. Geophys. Res., 101, 1943–1960, 1996.

Newell, R. E., Thouret, V., Cho, J. Y. N., Stoller, P., Marenco, A.,and Smit, H. G.: Ubiquity of quasi-horizontal layers in the tro-posphere, Nature, 398, 316–319, 1999.

Orville, R. E., Zisper, E. J., Brook, M., Weidman, C., Aulich, G.,Krider, E. P., Christian, H., Goodman, S., Blakeslee, R., andCummins, K.: Lightning in the region of the TOGA COARE,Bull. Am. Met. Soc., 78, 1055–1067, 1997.

Petersen, W. A., Rutledge, S. A., and Orville, R. E.: Cloud-to-ground lightning observations from TOGA COARE: Selected re-sults and lightning location algorithms, Mon. Weather Rev., 124,602–620, 1996.

Phillips, V. T. J. and Donner, L. J.: Cloud microphysics, radiationand vertical velocities in two- and three-dimensional simulationsof deep convection, Q. J. R. Meteorol. Soc., 132, 3011–3033,2006.

Pickering, K. E., Thompson, A. M., Dickerson, R. R., Luke, W. T.,McNamara, D. P., Greenberg, J. P., and Zimmerman, P. R.:Model calculations of tropospheric ozone production potentialfollowing observed convective events, J. Geophys. Res., 95,14 049–14 062, 1990.

Pickering, K. E., Thompson, A. M., Tao, W.-K., and Kucsera, T. L.:Upper tropospheric ozone production following mesoscale con-vection during STEP/EMEX, J. Geophys. Res., 98, 8737–8749,1993.

Pickering, K. E., Thompson, A. M., Wang, Y., Tao, W.-K., Mc-Namara, D. P., Kirchhoff, V. W. J. H., Heikes, B. G., Sachse,G. W., Bradshaw, J. D., Gregory, G. L., and Blake, D. R.: Con-vective transport of biomass burning emissions over Brazil dur-ing TRACE A, J. Geophys. Res., 101, 23 993–24 012, 1996.

Pickering, K. E., Wang, Y., Tao, W.-K., Price, C., and Muller, J.-F.: Vertical distributions of lightning NOx for use in regional andglobal chemical transport models, J. Geophys. Res., 103, 31 203–31 216, 1998.

Pickering, K. E., Thompson, A. M., Kim, H., DeCaria, A. J., Pfis-ter, L., Kucsera, T. L., Witte, J. C., Avery, M. A., Blake, D. R.,Crawford, J. H., Heikes, B. G., Sachse, G. W., Sandholm, S. T.,and Talbot, R. W.: Trace gas transport and scavenging in PEM-Tropics B South Pacific Convergence Zone convection, J. Geo-phys. Res., 106, 32 591–32 602, 2001.



Poulida, O., Dickerson, R. R., and Heymsfield, A.: Stratosphere-troposphere exchange in a midlatitude mesoscale convectivecomplex, J. Geophys. Res., 101, 6823–6836, 1996.

Price, C. and Rind, D.: A simple lightning parameterization forcalculating global lightning distributions, J. Geophys. Res., 97,9919–9933, 1992.

Price, C. and Rind, D.: What determines the cloud-to-ground light-ning fraction in thunderstorms?, Geophys. Res. Lett., 20, 463–466, 1993.

Price, C., Penner, J., and Prather, M.: NOx from lightning: 1. Globaldistribution based on lightning physics, J. Geophys. Res., 102,5929–5941, 1997.

Ramanathan, V. R., Dirks, R., Grossman, R., Heymsfield, A., Kuet-tner, J., and Valero, F. P. J.: Central Equatorial Pacific design,Center for Clouds, Chemistry, and Climate, University of Cali-fornia, San Diego, 1993.

Ray, P. S., MacGorman, D. R., Rust, W. D., Taylor W. L., and Wal-ters Rasmussen, L.: Lightning location relative to storm structurein a supercell and a multicell storm, J. Geophys. Res., 92, 5713–5724, 1987.

Redelsperger, J., Brown, P. R. A., Guichard, F., Hoff, C., Kawasima,M., Lang, S., Montmerle, T., Nakamura, K., Saito, K., Seman,C., Tao, W.-K., and Donner, L. J.: A GCSS model intercompari-son for a tropical squall line observed during TOGA-COARE. I:Cloud-resolving models, Q. J. R. Meteorol. Soc., 26, 823–863,2000.

Salzmann, M.: Influences of deep convective cloud systemson tropospheric trace gases and photochemistry over thetropical West Pacific: A modeling case study, Ph.D. thesis,Johannes Gutenberg-Universitat Mainz, Mainz, Germany,http://nbn-resolving.de/urn/resolver.pl?urn=urn:nbn:de:hebis:77-9470, 2005.

Salzmann, M., Lawrence, M. G., Phillips, V. T. J., and Donner,L. J.: Modelling tracer transport by a cumulus ensemble: Lat-eral boundary conditions and large-scale ascent, Atmos. Chem.Phys., 4, 1797–1811, 2004,http://www.atmos-chem-phys.net/4/1797/2004/.

Salzmann, M., Lawrence, M. G., Phillips, V. T. J., and Donner, L. J.:Model sensitivity studies regarding the role of the retention co-efficient for the scavenging and redistribution of highly solubletrace gases by deep convective cloud systems., Atmos. Chem.Phys., 7, 2027–2045, 2007.

Sander, R. and Crutzen, P. J.: Model study indicating halogen ac-tivation and ozone destruction in polluted air masses transportedto the sea, J. Geophys. Res., 101, 9121–9138, 1996.

Scala, J. R., Garstang, M., Tao, W.-K., Pickering, K. E., Thomp-son, A. M., Simpson, J., Kirchhoff, V. W. J. H., Browell, E. V.,Sachse, G. W., Torres, A. L., Gregory, G. L., Rasmussen, R. A.,and Khalil, M. A. K.: Cloud draft structures and trace gas trans-port, J. Geophys. Res., 95, 17 015–17 030, 1990.

Schumann, U. and Huntrieser, H.: The global lightning-induced ni-trogen oxides source, Atmos. Chem. Phys., 7, 3823–3907, 2007,http://www.atmos-chem-phys.net/7/3823/2007/.

Skamarock, W. C., Klemp, J. B., and Dudhia, J.: Prototypes for theWRF (Weather Research and Forecasting) model, in: Preprints,Ninth Conf. Mesoscale Processes, pp. J11–J15, Amer. Meteor.Soc., Fort Lauderdale, FL, 2001.

Skamarock, W. C., Dye, J. E., Barth, M. C., Stith, J. L., Ridley,B. A., and Baumann, K.: Observational- and modeling-basedbudget of lightning-produced NOx in a continental thunderstorm,J. Geophys. Res., 108, 4035, doi:10.1029/2002JD002163, 2003.

Stenchikov, G., Dickerson, R., Pickering, K., Ellis Jr., W., Dod-dridge, B., Kondragunta, S., Poulida, O., Scala, J., and Tao,W.-K.: Stratosphere-troposphere exchange in a midlatitudemesoscale convective complex, 2. Numerical simulations, J.Geophys. Res., 101, 6837–6851, 1996.

Takemi, T. and Rotunno, R.: The effects of subgrid model mixingand numerical filtering in simulations of mesoscale cloud sys-tems, Mon. Weather Rev., 131, 2085–2191, 2003.

von Kuhlmann, R., Lawrence, M. G., Crutzen, P. J., and Rasch,P. J.: A model for studies of tropospheric ozone and nonmethanehydrocarbons: Model description and ozone results, J. Geophys.Res., 108, doi:10.1029/2002JD002893, 2003.

Walcek, C. J.: Minor flux adjustment near mixing ratio extremesfor simplified yet highly accurate monotonic calculation of traceradvection, J. Geophys. Res., 105, 9335–9348, 2000.

Wang, C. and Chang, J. S.: A three-dimensional numerical model ofcloud dynamics, microphysics, and chemistry: 3. Redistributionof pollutants, J. Geophys. Res., 98, 16 787–16 798, 1993.

Wang, C. and Prinn, R. G.: On the roles of deep convective cloudsin tropospheric chemistry, J. Geophys. Res., 105, 22 269–22 297,2000.

Wang, C., Crutzen, P. J., Ramanathan, V., and Williams, S. F.: Therole of a deep convective storm over the tropical Pacific Ocean inthe redistribution of atmospheric chemical species, J. Geophys.Res., 100, 11 509–11 516, 1995.

Webster, P. J. and Lukas, R.: TOGA COARE: The Coupled Ocean-Atmosphere Response Experiment, Bull. Am. Met. Soc., 73,1377–1416, 1992.

Wild, O.: Modelling the global tropospheric ozone budget: explor-ing the variability in curerent models, Atmos. Chem. Phys., 7,2643–2660, 2007,http://www.atmos-chem-phys.net/7/2643/2007/.

Yanai, M., Chen, B., and Tung, W.-W.: The Madden-Julian Oscil-lation observed during the TOGA COARE IOP: Global view, J.Atmos. Sci., 57, 2374–2396, 2000.


http://nbn-resolving.de/urn/resolver.pl?urn=urn:nbn:de:hebis:77-9470

http://nbn-resolving.de/urn/resolver.pl?urn=urn:nbn:de:hebis:77-9470

http://www.atmos-chem-phys.net/4/1797/2004/



cloud system resolving model study of the roles of deep ...gated in a 7-day 3-d 248×248km2...

Documents